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13

Volcanism, the atmosphere and climate through time

anja schmidt and alan robock

13.1 Introduction

Volcanism can affect the environment on timescales of weather (days to weeks)and climate (months to years). In the past 250 years, the atmospheric andclimatic effects of several volcanic eruptions have been witnessed and docu-mented by scientist and non-scientist alike. For example, the 1783–1784 Lakieruption in Iceland, which was followed by hot summer temperatures andcold winter temperatures in central Europe (Thordarson and Self, 2003; Omanet al., 2006a; Schmidt et al., 2012b), was described in great detail by an Icelandicpriest (Steingrímsson, 1788). The 1815 Tambora eruption in Indonesia causedthe ‘Year Without a Summer’ in 1816, which inspired Mary Shelley’sFrankenstein.

The scientific understanding of the climatic effects caused by short-lived explo-sive volcanic eruptions has advanced greatly in the past five decades, mainly dueto theoretical work and observations following the eruptions of Agung in 1963,Mount St Helens in 1980, El Chichón in 1982 and, in particular, Mount Pinatuboin 1991 (for reviews see Robock, 2000; Timmreck, 2012).

In the past two decades, the climate relevance of effusive volcanic activityinjecting gases mainly into the troposphere has become more recognized (e.g. Grafet al., 1997; Schmidt et al., 2012a), and also the atmospheric and climatic effectsof small- to moderate-sized explosive eruptions (e.g. Solomon et al., 2011; Neelyet al., 2013). However, we are only beginning to understand how large-volumeflood basalt eruptions such as the Deccan Traps in India at around 65 Ma mayhave affected the environment.

This chapter provides an overview of the impact of various eruption types andstyles on Earth’s atmosphere and climate through time.

Volcanism and Global Environmental Change, eds. Anja Schmidt, Kirsten E. Fristad and Linda T. Elkins-Tanton.Published by Cambridge University Press. © Cambridge University Press 2015.

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13.2 Volcanism on Earth

The size of a volcanic eruption can be expressed in terms of the volume (km3)or mass (kg) of magma produced. The Volcanic Explosivity Index (VEI)uses volcanological data such as eruption column height and ejecta volume tomeasure the relative magnitude of an eruption’s explosivity using an open-endedlogarithmic scale (except for VEI between 0 and 2) (Newhall and Self, 1982).The fragmentation of volatile-rich magma during explosive volcanic eruptionsproduces large quantities of volcanic tephra, and generally the likelihood thatejecta reach stratospheric altitudes increases with increasing VEI. There are,however, examples of VEI 5 eruptions, such as the 1980 Mount St Helenseruption, which injected little sulfur into the stratosphere, and thus had no impacton climate (Robock, 1981).

One of the largest eruptions in the twentieth century, the 1991 Mount Pinatuboeruption, had a VEI of 6 and produced a bulk magma volume of about 10 km3.For context, the 1815 Tambora eruption (VEI 7) had a bulk volume > 100 km3.The 1783–1784 Laki eruption in Iceland greatly exceeded twentieth-century lavavolumes, yielding about 15 km3 of extrusives in eight months (Thordarson andSelf, 2003).

Past episodes of continental flood basalt (CFB) volcanism produced hugelava volumes on the order of tens of millions of cubic kilometres. Examplesinclude the emplacement of the Deccan Traps, 65 Ma (Late Cretaceous;> 106 km3); or the Columbia River Basalt Group (Western United States),17–10 Ma (Middle Miocene; ~ 210,000 km3), including the Roza flow,14.7 Ma (~ 1,300 km3). As also discussed in Chapters 5 and 11, CFB volcanismis typified by numerous, recurring large-volume eruptive phases, each lastingdecades or longer. On geological timescales, the climactic pulses that buildup an entire CFB province are short-lived – typically about 1 Ma or less. Forthe assessment of the environmental perturbations it is, however, importantto recognize the pulsed nature of CFB volcanism, with the duration of non-eruptive phases (usually centuries to millennia) outlasting the duration ofindividual eruptive phases (years to decades). An individual decade-longeruptive phase would have yielded between 103 and 104 km3 of extrusivesbuilding up pāhoehoe-dominated lava flow fields (see Chapter 11 for details;Self et al., 2006; 2008). Despite emplacing huge magma volumes, individualeruptions would rarely have exceeded a VEI of 4 and, typically, volcanic gaseswould have been injected into the upper troposphere and lower stratosphere(see Chapter 11). There is, however, evidence that some CFB provinces wereassociated with more violent silicic eruptions (e.g. Bryan and Ferrari, 2013; seealso Chapter 1].

196 Schmidt and Robock

13.3 Volcanic gas emissions

Water vapour (H2O) and carbon dioxide (CO2) are the most abundant volatilespecies released during a volcanic eruption, but in the short term their effect onatmospheric composition is negligible because of the insignificant relative contri-bution to the high atmospheric background concentrations of H2O and CO2.However, on the timescale of the age of Earth, volcanic ‘outgassing’ has beenthe source of our current atmosphere (see also Chapter 14).

Compared to 35,000 Teragrams (Tg) of anthropogenic CO2 emissions per year,estimates of the CO2 flux from present-day subaerial and submarine volcanismrange between 130 and 440 Tg (Gerlach, 2011, and references therein). Self et al.(2008) estimated that the Deccan Traps released up to 14 Tg of CO2 per km

3 oflava erupted (assuming a degassing efficiency of 80%). For a typical decade-longCFB eruptive phase producing a total lava volume of 1,000 km3 this equatesto 1400 Tg of CO2 per year, which is one order of magnitude smaller than thecurrent anthropogenic flux.

Volcanic eruptions also release halogen species (mainly bromine oxide, hydro-gen chloride, hydrogen bromide and hydrogen fluoride), which play an importantrole in volcanic plume chemistry (von Glasow, 2010). Over the industrialera, anthropogenic chlorofluorocarbon (CFC) emissions have resulted in highstratospheric chlorine and bromine concentrations. These halogens destroy strato-spheric O3 efficiently, leading to phenomena such as the Antarctic ‘ozone hole’.Stratospheric O3 depletion was also observed after the eruptions of El Chichónin 1982 and Mount Pinatubo in 1991 because volcanic aerosol particles serveas surfaces for heterogeneous reactions promoting conversion of less reactive(anthropogenic) chlorine/bromine species into more reactive forms (Solomonet al., 1998; Solomon, 1999 for a review). Therefore, in general, the chemicalimpact of volcanic eruptions is to exacerbate anthropogenic-driven O3 destruction.Anthropogenic chlorine concentrations are steadily declining; hence, future erup-tions will not deplete stratospheric O3 except if an eruption itself injects sufficientchlorine (in the form of hydrochloric acid) into the stratosphere; this is a matterof debate (e.g. Tabazadeh and Turco, 1993; Kutterolf et al., 2013; see alsoChapter 16). Chapter 20 discusses how halogen-bearing species emitted by floodbasalt eruptions could have resulted in global-scale O3 depletion during the end-Permian (252 Ma) emplacement of the Siberian Trap province.

To date, sulfur dioxide (SO2) is the sole volcanic volatile species that has beenobserved to alter the radiative balance of the atmosphere, because of its conversionto volcanic sulfuric acid aerosol. Sulfur species contribute between 2% and 35% byvolume of the gas phase, and SO2 and hydrogen sulfide (H2S) are most abundant.In the troposphere, H2S rapidly oxidizes to SO2, which commonly has a chemical

Volcanism, the atmosphere, climate through time 197

lifetime of hours to days. If released into the stratosphere, the lifetime of SO2

increases to about 3 weeks due to slower removal processes than in thetroposphere. SO2 is removed from the atmosphere via dry and wet deposition(‘acid rain/snow’) or oxidization to form sulfuric acid aerosol. Gas-phase oxidationof SO2 by the hydroxyl radical (OH�) forms sulfuric acid (H2SO4) vapour, which isa low volatility compound that forms new particles (‘nucleation’) and/or rapidlycondenses to grow existing particles to larger sizes. Within clouds, SO2 undergoesaqueous-phase oxidation via reactions with dissolved hydrogen peroxide (H2O2)or ozone (O3). Volcanic sulfuric acid particles are typically composed of 75%by weight of H2SO4 and 25% by weight of H2O (Bekki, 1995).

There has not been a large stratospheric injection of volcanic SO2 sincethe 1991 Mount Pinatubo eruption, which released about 20 Tg of SO2 into thestratosphere. For comparison, the current anthropogenic SO2 flux is about 116 Tgper year. The 1982 El Chichón eruption injected ~ 7 Tg of SO2 into the strato-sphere. Between the years 2000 and 2012 there has been a series of small- tomoderate-sized eruptions (VEI 3–4) injecting at the most ~ 1.3 Tg of SO2 duringan individual eruption (Nabro, 2011, in Eritrea). In comparison, individual decade-long CFB eruptive phases would have released up to 5,000 Tg of SO2 (Self et al.,2008). The Roza flow, the best-studied CFB eruptive phase, released about1,200 Tg of SO2 per year for a decade or longer into altitudes between 7 and13 km (Thordarson and Self, 1996). For context, Laki was one of the largesthistoric flood basalt events and injected about 120 Tg of SO2 in 8 months(Thordarson and Self, 2003, and references therein).

The annual sulfur flux from continuously degassing and sporadically eruptingvolcanoes is about 13 Tg, based on measurements at 49 volcanoes ‘only’ (Andresand Kasgnoc, 1998). The uncertainties on this estimate are large (about � 50%);hence, there is also a large uncertainty on the magnitude of the radiative effectsinduced by these emissions (Graf et al., 1997; Schmidt et al., 2012a). Withthe advent of the satellite era, more numerous measurements of the volcanic sulfurflux have become available (e.g. Carn et al., 2013) but the picture is still farfrom complete.

13.4 Atmospheric and climatic impacts of volcanic aerosol

Figure 13.1 illustrates how volcanic gases and particles affect atmospheric com-position, chemistry and Earth’s climate. Tephra is a major constituent of a volcaniccloud associated with explosive eruptions. Airborne tephra particles rapidly fall outof the atmosphere within minutes to days, hence their climatic effects are mostlynegligible (Robock, 1981; Niemeier et al., 2009). Continent-sized ash depositsproduced by super-eruptions (VEI � 8), however, may induce short-term climate


change (Jones et al., 2007), and Chapter 17 discusses the environmental impactsassociated with ash deposition.

In contrast to airborne volcanic tephra, volcanic sulfuric acid aerosol particlescan alter the radiative balance of Earth on timescales relevant for climate changedue to their ability to scatter and absorb solar radiation (aerosol direct effect).Sulfuric acid aerosol scatters radiation across the entire solar spectrum due to itsoptical properties and typical particle radius of 0.5 μm (although there is a smalldegree of absorption in the near-infrared spectrum). Once in the stratosphere,volcanic sulfuric acid aerosol has a typical e-folding time of 9–12 monthsfor tropical eruptions and 2–4 months for high-latitude eruptions (Kravitz andRobock, 2011). The top of a stratospheric aerosol cloud absorbs solar radiation inthe near-infrared, whereas the lower stratosphere is heated by absorption of upwardthermal infrared radiation (Stenchikov et al., 1998). An aerosol cloud also forwardscatters some of the incoming solar radiation, resulting in enhanced downwarddiffuse radiation, thereby increasing the sky’s brightness. The net radiative effectof explosive volcanic eruptions is, however, a reduction in surface temperaturesbecause the scattering exceeds the absorption efficiency.

In the case of the 1991 eruption of Mount Pinatubo, stratospheric aerosolconcentrations remained elevated for about 2 years, and globally averaged strato-spheric temperatures increased by around 2–3 K 3 months after the eruption

Figure 13.1. Schematic diagram showing the effect of volcanic gases andparticles on the atmospheric composition and Earth’s climate. Adapted from Plate1 of Robock (2000).


(Labitzke and McCormick, 1992). About 18 months following the June1991 Mount Pinatubo eruption, the peak reduction in globally averaged lower-tropospheric temperatures reached 0.5 K, with climate models predicting acooling of the same magnitude when including water-vapour feedbacks andremoving the El Niño–Southern Oscillation signal (Soden et al., 2002, andreferences therein).

Kirchner et al. (1999) showed that the aerosol formed after tropical eruptionsresults in an enhanced pole-to-equator heating gradient in the lower stratospherethat creates a stronger polar vortex and associated positive mode of the ArcticOscillation in tropospheric circulation. This results in a warming over northernAmerica, northern Europe and Russia, and a cooling over the Middle East, duringthe first winter, and sometimes the second winter, following a tropical eruption(Robock and Mao, 1992; Graf et al., 1993). This indirect advective effect ontemperature is stronger than the direct radiative effect that dominates at lowerlatitudes and during summer. Figure 13.2 highlights that while observationsconfirm this ‘winter warming’ response the majority of current climate modelsfail to reproduce the magnitude of this effect (Driscoll et al., 2012).

Research on tropical and high-latitude eruptions revealed that feedbacksbetween the additional aerosol loading and atmospheric dynamics can weakenthe Asian and African summer monsoon systems (Oman et al., 2006b). For high-latitude eruptions in particular, the season of eruption is important for assessingthe magnitude of the climatic response (Schmidt et al., 2010; Kravitz andRobock, 2011). Trenberth and Dai (2007) showed that after the 1991 MountPinatubo eruption precipitation over land and continental freshwater dischargedecreased significantly between October 1991 and September 1992. FollowingRobock and Liu (1994), Haywood et al. (2013) showed that the asymmetricaerosol loading between the northern and southern hemispheres after eruptionscan influence the sea surface temperature gradient in the Atlantic, which in turnaffects Sahelian precipitation rates (three of the four driest Sahelian summersbetween 1900 to 2010 were preceded by an eruption in the northernhemisphere).

Figure 13.1 also shows that aside from the aerosol direct effect, aerosol particlescan also alter cloud amount and albedo by acting as cloud condensation nucleior ice nuclei, thus modifying the optical properties and lifetime of clouds (‘aerosolindirect effects’). Whether volcanic particles alter cirrus cloud properties has beenthe subject of several studies with contradicting results (e.g. Sassen, 1992;Luo et al., 2002; see Robock et al., 2013 for a review). In contrast, the abilityof volcanic sulfates to act as cloud condensation nuclei and their indirecteffect on low-level warm clouds has been confirmed by means of measurements


(e.g. Mather et al., 2003), satellite retrievals (e.g. Gassó, 2008), and modelling(e.g. Graf et al., 1997; Schmidt et al., 2010, 2012a).

13.5 Volcano-climate interactions through time

Rampino and Self (1982) noted that eruption size does not correlate with themagnitude of the climatic effect. Pinto et al. (1989) used a one-dimensionalmicrophysics model to demonstrate that the volcanic impact on climate becomesself-limited with increasing SO2 release because microphysical processes such ascoagulation cause particles to grow to large sizes, which have a lower optical depthper unit mass and fall out of the stratosphere faster. Timmreck et al. (2010) andEnglish et al. (2013) simulated an eruption thought to be representative of theYoungest Toba Tuff eruption (74 ka) using three-dimensional microphysics

Figure 13.2 Comparison of the surface temperature changes (in K) as observed(a) and simulated by current climate models (b) averaged for two winter seasonsafter a major eruption in the tropics. The difference between modelled andobserved surface temperature changes is shown in (c) with dashed contours (bluecolours in colour version) indicating that modelled temperatures are too cold andgrey-shaded contours (yellow colours in colour version) indicating that modelledtemperatures are too warm compared to the observations. Figures reproducedbased on data analysed and presented in Driscoll et al. (2012). A black and whiteversion of this figure will appear in some formats. For the colour version, pleaserefer to the plate section.


models. These model simulations suggested a global mean temperature responseabout three times weaker (change of about –3.5 K) than previously estimatedusing climate models that did not consider microphysical processes. Pinto et al.(1989) and Bekki (1995) noted that oxidant depletion of OH� could prolongthe climate response producing sulfuric acid aerosol more slowly and increasingthe lifetime of SO2, but Robock et al. (2009) suggested that this would be asmall effect.

A reasonable rule of thumb might be that a minimum amount of 5 Tg of SO2

injected into the stratosphere is likely to significantly alter Earth’s radiativebalance on timescales of months to years; however, there are complicatingfactors such as the season and latitude of an eruption as well as what temperaturechange we consider significant. In any case, simply extrapolating the magnitudeof the climatic effects of present-day explosive volcanism to, for example,CFB volcanism is bound to be flawed due to the non-linear relationshipbetween eruption magnitude and climatic effect, as well as the differences ineruption style.

13.5.1 Ancient large-scale CFB eruptions

Five Phanerozoic periods of CFB volcanism in particular have been the subject ofa long-standing debate about their association with major environmental changesas evident from the proxy record of, for example, the abundances and diversity ofplanktonic foraminifera. However, the mechanisms by which CFB volcanismmay have triggered these environmental changes remain enigmatic (for reviewssee Officer et al., 1987 and Wignall, 2001; see also Chapter 11).

Most studies on CFB volcanism suggested, by analogy to present-day volcan-ism, a short-term cooling effect (lasting years to decades) from SO2 and theformation of sulfuric acid aerosol; together with a long-term warming effect(lasting tens to thousands of years) from the seemingly ‘high’ volcanic CO2

emissions. However, the annual volcanic CO2 flux during an eruptive phaseequates to about 4% of the current annual anthropogenic flux (Section 13.3).To date, there has been only one carbon-cycle modelling study, and it found anegligible effect of volcanic CO2 emissions from the Deccan Traps on LateCretaceous temperatures (Caldeira and Rampino, 1990). Whether volcanic CO2

affected long-term weathering rates and/or ocean acidity could be addressed infuture using carbon-cycle models. Alternative explanations for the long-termwarming trend observed at the time of, for example, the Siberian Traps revolvearound the dissociation of gas hydrates (see also Chapters 10, 12 and 20).Release of isotopically light carbon from gas hydrates may explain the negative


carbon isotope excursions at the end-Permian, rather than these excursions beingdue to the release of volcanic carbon (Wignall, 2001, and references therein).

For the 14.7 Ma decade-long Roza flow, Thordarson and Self (1996) estimated achange in aerosol optical depth (AOD) at 550 nm of between 7 and 13, which is atleast 2.5 times larger than that for Toba (peak AOD of 2.6; English et al., 2013).However, Thordarson and Self (1996) assumed that all SO2 released is convertedto sulfuric acid aerosol of a certain size. Accounting for oxidant availability andmicrophysical processes in a global model simulation of a Roza-scale eruption,A. Schmidt (unpublished data) found an annual mean AOD change of 1.2 whenaveraged over the northern hemisphere (Figure 13.3), which is significantlylower than previously estimated but nonetheless would have resulted in asubstantial perturbation of climate if maintained for a decade. Since the climaticperturbations are relatively short-lived on geological timescales, A. Schmidt(unpublished data) suggested that better constraints on the frequency and dur-ation of individual eruptions as well as non-eruptive phases are needed to fullyquantify the magnitude and duration of the climatic and environmental effects(see also Chapter 11).

13.5.2 Post-2000 small- to moderate-sized eruptions (VEI 3–4)

The rate of global warming slowed between the years 2005 and 2012. Among othercontributing factors, VEI 3–4 volcanic eruptions during that period contributed a

(a) (b)

Figure 13.3 Aerosol–climate model simulation of a decade-long Roza-scaleeruption (releasing 1200 Tg of SO2 per year) with (a) showing the annual meanchange in aerosol optical depth (AOD) at 550 nm for the tenth year of the eruptionwith respect to a pre-industrial control simulation, and (b) showing the latitudinalmean AOD for a pre-industrial control simulation (dashed) and the volcanicallyperturbed simulation (solid). The latitude of the eruption is indicated by the blacktriangle.


small negative radiative forcing by increasing stratospheric aerosol concentrations(Solomon et al., 2011; Vernier et al., 2011; Neely et al., 2013). Fyfe et al. (2013)estimated a global surface temperature response of at least –0.07�C.

The June 2011 eruption of Nabro (VEI 3–4) injected about 1.3 Tg of SO2 intothe upper troposphere (9–14 km), where seemingly the climate impacts would havebeen more muted than for an injection into the stratosphere. Bourassa et al. (2012),however, suggested that deep convection and interaction with the Asian summermonsoon lofted Nabro’s SO2 into the stratosphere, hence prolonging the aerosolradiative forcing.

These recent findings highlight the necessity for detailed observations ofsmall-magnitude eruptions (VEI of 3 to 4) to allow for a better understandingof present-day climate change. Such findings might also change our view of whateruption magnitude and mass of SO2 released into the atmosphere we needto consider relevant for climate change on decadal timescales, and also provideinformation of relevance for stratospheric geoengineering proposals with sulfuricacid aerosol.

13.6 Summary

Observations have confirmed that volcanic eruptions can have numerous impactson our environment, mainly via the sulfuric acid aerosol direct effects, resultingin a reduction of surface temperatures, and changes in water-vapour concen-trations and stratospheric heating rates. The latter results in pole-to-equatorgradients that induce an indirect advective effect on surface temperatures referredto as winter warming. Transient and asymmetric volcanic aerosol forcingshave been shown to weaken the Asian and African summer monsoons and tocause precipitation deficits in the Sahel. There is no observational evidencethat eruptions force certain circulation modes such as El Niño but the issueremains a matter of debate. There is increasing evidence for an aerosol indirecteffect on liquid water clouds induced by volcanic sulfuric acid particles butwhether, and to what extent, explosive volcanic eruptions affect ice cloudsremains debated.

Future eruptions will present an opportunity to test our current understan-ding and fill gaps in our ability to monitor volcanic clouds and to obtaindetailed measurements. Climate models and observations have greatlyenhanced our understanding of the climatic impacts of eruptions. However,we have not yet obtained sufficient observations of, for example, the temporalevolution of the aerosol size distribution in both the stratosphere andtroposphere for different eruption sizes, which will be an acid test for climateand aerosol models.


Acknowledgements

We thank Steve Self for a review that helped to improve this chapter. Anja Schmidtis funded by an Academic Fellowship from the School of Earth and Environment,University of Leeds. Alan Robock is supported by US National Science Founda-tion grants AGS-1157525 and CBET-1240507. We are very grateful to SimonDriscoll for providing the data for Figure 13.2.

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